Differential uplift along the northern margin of the central anatolian plateau: Inferences from marine terraces

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Differential uplift along the northern margin of the Central Anatolian

Plateau: inferences from marine terraces

Cengiz Yildirim

a,b,c,*

_{, Daniel Melnick}

c

_{, Paolo Ballato}

c

_{, Taylor F. Schildgen}

c

_,

Helmut Echtler

b

, A. Evren Erginal

d

, Na

_{ﬁye Güneç Kıyak}

e

, Manfred R. Strecker

c

a_Istanbul Technical University, Eurasia Institute of Earth Sciences, Turkey

b_{Deutsches GeoForschungsZentrum (GFZ), Potsdam, Germany}

c_{Institut für Erd- und Umweltwissenschaften, and DFG Leibniz Center for Surface, Process and Climate Studies, Universität Potsdam, Potsdam, Germany}

d_{Department of Geography, Ardahan University, Turkey}

e_{Department of Physics, Is¸ık University, _Istanbul, Turkey}

a r t i c l e i n f o

Article history: Received 26 March 2013 Received in revised form 29 August 2013

Accepted 13 September 2013 Available online 10 October 2013 Keywords:

Orogenic Plateaus Central Anatolian Plateau Plateau margins Central Pontides Orogenic wedges Black Sea Marine terraces Uplift rate

a b s t r a c t

Emerged marine terraces and paleoshorelines along plate margins are prominent geomorphic markers that can be used to quantify the rates and patterns of crustal deformation. The northern margin of the Central Anatolian Plateau has been interpreted as an actively deforming orogenic wedge between the North Anatolian Fault and the Black Sea. Here we use uplifted marine terraces across principal faults on the Sinop Peninsula at the central northern side of the Pontide orogenic wedge to unravel patterns of Quaternary faulting and orogenic wedge behavior. We leveled the present-day elevations of paleo-shorelines and dated marine terrace deposits using optically stimulated luminescence (OSL) to determine coastal uplift. The elevations of the paleoshorelines vary between 4 0.2 and 67 1.4 m above sea level and OSL ages suggest terrace formation episodes during interglacial periods at ca 125, 190, 400 and 570 ka, corresponding to marine isotopic stages (MIS) 5e, 7a, 11 and 15. Mean apparent vertical displacement rates (without eustatic correction) deduced from these terraces range between 0.02 and 0.18 mm/a, with intermittent faster rates of up to 0.26 mm/a. We obtained higher rates at the eastern and southern parts of the peninsula, toward the hinterland, indicating non-uniform uplift across the different morphotectonic segments of the peninsula. Our data are consistent with active on- and offshore faulting across the Sinop Peninsula. When integrated with regional tectonic observations, the faulting pattern reﬂects shortening distributed over a broad region of the northern margin of the Central Anatolian Plateau during the Quaternary.

1. Introduction

Along coastal margins, emerged marine terraces form tectonic bench markers that record the rates and patterns of crustal defor-mation (e.g., Bloom et al., 1974; Merritts and Bull, 1989; Pedoja et al., 2006; Bishop, 2007; Caputo et al., 2010; Melnick et al., 2012a). Emerged marine terraces in these environments are created by a combination of sea-level oscillations and coeval ver-tical crustal motions (e.g.,Lajoie, 1986). The total displacement of a marine terrace is the averaged sum of coseismic and interseismic movements with respect to the position of the corresponding paleo sea-level during terrace formation, and is therefore commonly

analyzed to deduce uplift rates over intermediate timescales (e.g.,

Lajoie, 1986). As such, the spatio-temporal variation of uplift rates recorded by marine terraces can provide information about pat-terns of deformation from the scale of individual faults (e.g.,

Valensise and Ward, 1991; Caputo et al., 2010; Melnick et al., 2012b) to the scale of tectonic plate motions (Gardner et al., 1992; Westaway, 1993; Armijo et al., 1996; Pedoja et al., 2011).

The margins of the Central Anatolian orogenic plateau are delimited by the Central Tauride mountains in the south, which border the Mediterranean Sea, and the Central Pontide mountains in the north, which abut the Black Sea (Fig. 1A). A segment of the shoreline of the Black Sea was coseismically uplifted between 0.5 and 0.7 m (Wedding, 1969; Ketin and Abdüsselamoglu, 1970) dur-ing the Ms 6.6 Bartın earthquake in 1968 as a result of offshore thrust faulting (McKenzie, 1972) associated with ongoing short-ening along the western part of the northern plateau margin

* Corresponding author. _Istanbul Technical University, Eurasia Institute of Earth Sciences, Turkey.

E-mail address:cyildirim@itu.edu.tr(C. Yildirim).

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(Fig. 1B). With respect to longer timescales, however, the defor-mation and surface-uplift history of the Central Pontides is not very well known and subject to ongoing debate (Erinç and _Inandık, 1955; Akkan, 1975; Karabıyıkoglu, 1984; Yildirim et al., 2011).

Yildirim et al. (2011)suggested that the broad restraining bend of the North Anatolian Fault (NAF) has led to the development of an active orogenic wedge (e.g., Dahlen, 1990) that drives crustal thickening and uplift of the Ilgaz and Sinop ranges (Fig. 1C). Given that the NAF is inferred to have initiated in the late Miocene to early Pliocene (e.g.,Hubert-Ferrari et al., 2002), such a scenario would

imply shortening and uplift since that time.Mazzini et al. (2013)

also found that the stable isotope composition of soil carbonates ranging in age from MN9 (ca 11 Ma) to MN11 (ca 8 Ma) inland of the Central Pontides are less depleted compared to that of modern precipitation values (Schemmel et al., 2013), indicating that the Central Pontides have become a larger barrier to precipitation since the late Miocene.

The Sinop Peninsula constitutes the northernmost promontory of the Central Pontides (Fig. 1A, B). The peninsula preserves several ﬂights of marine terraces that have emerged above sea level (Erinç

Fig. 1. (A). Simpliﬁed neotectonic map of Turkey (afterBarka and Reilinger, 1997;S¸engör et al., 1985; Schildgen et al., 2012). CAP: Central Anatolian Plateau; EAF: East Anatolian

Fault; NAF: North Anatolian Fault, WAEP: Western Anatolian Extensional Province; BZSZ: Bitlis-Zagros Suture Zone. (B) Active faults in the study area and focal mechanism solutions of major earthquakes (McKenzie, 1972; Finetti et al., 1988; Dirik, 1993; Andrieux et al., 1995; Barka and Reilinger, 1997; Reilinger et al., 2006). Years are labeled on epicenters of the 1943 Tosya, 1944 Bolu-Gerede, 1951 Kurs¸unlu and 1968 Bartın earthquakes. Moment magnitudes (M) of other earthquakes are indicated with colored symbols. Bathymetric contours (white lines, 250-m interval) were extracted from GTOPO30 data (http://edcwww.cr.usgs.gov/landdaac/gtopo30/gtopo30html). BF: Balıfakı Fault; CF: Cide Faults; EF: Erikli Fault; EkF: Ekinveren Fault; KF: Karabük Fault; ILR: Ilgaz Range; TB: Tosya Basin. (C) Simpliﬁed conceptual model of the northern margin of the Central Anatolian Plateau (modiﬁed

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and _Inandık, 1955; Akkan, 1975; Karabıyıkoglu, 1984; Yildirim et al., 2011). In this study, we focus on marine terraces located along different morphotectonic segments of the Sinop Peninsula to evaluate the rates and patterns of onshore and offshore deforma-tion. We present (1) the geomorphic characteristics of uplifted terraces including precise shoreline-angle measurements; (2) ages of the terrace deposits based on optically-stimulated luminescence (OSL) measurements and their correlation with marine isotopic substages (MIS); (3) the net vertical displacements of the marine terraces; (4) mean and intermediate uplift rates along the coastal zone of the northern Anatolian plateau margin; and (5) an

evaluation of on- and offshore structures in the context of the regional tectonic framework. The tectonic geomorphology of this region is well suited to assess deformation processes on interme-diate timescales, which may allow for a linkage of active defor-mation with long-term topographic development. Speciﬁcally, we test whether or not active offshore and onshore faults have accommodated signiﬁcant uplift on intermediate (Pleistocene) timescales. Taken together, this information can help to better constrain the spatiotemporal pattern of deformation and thus the seismotectonic segmentation of the northern margin of the Central Anatolia Plateau.

2. Tectonic and geologic setting

The Sinop Peninsula is one of the few places along the Turkish Black Sea coast where folded Miocene marine sediments and severalﬂights of emerged Quaternary marine abrasion platforms have been described (Erinç and _Inandık, 1955; _Inandık, 1955; Akkan, 1975; Karabıyıkoglu, 1984; Yildirim et al., 2011). The Black Sea is an oceanic basin more than 2000 m deep that developed due to back-arc extension associated with the northward subduction of the Tethyan plate during the Mesozoic (Okay et al., 1994; Okay and Tüysüz, 1999). Currently, the Turkish Black Sea coast is under compression due to the AfricaeArabia and Eurasia collision (Finetti et al., 1988; Barka and Reilinger, 1997; Cloetingh et al., 2003; Munteanu et al., 2011).

The structural, geomorphic and kinematic characteristics of the Sinop region suggest that the northern margin of the Central Anatolian Plateau constitutes an active north-vergent accretionary orogenic wedge, with strain broadly partitioned between the North Anatolian Fault and the abyssal plain of the Black Sea (Fig. 1C;

Yildirim et al., 2011). The Sinop Peninsula is a transitional zone be-tween the abyssal plain of the Black Sea and the array of high ranges of the northern margin (Fig. 1B). An offshore seismic reflection profile located north of the Sinop Peninsula shows back-stepping imbricate thrusts on a low-angle detachment that have deformed abyssal-plain sediments since the late Miocene (Finetti et al., 1988). Onshore, industry-style seismic reflection profiles across the Sinop Range and the Sinop Peninsula reveal thrust imbricates similar to the offshore structures (Ayd_{ın et al.,1995}). In contrast, using shallow seismic reflection profiles,Özhan (1989)identified normal faults affecting the continental shelf between these two major shortening structures. From the seismic re_{flection profile of}Özhan (1989)it is difficult to estimate the depth extent of these normal faults, but they most likely only affect the uppermost crustal levels (200e250 ms), as they are not observed on deeper profiles.

The Balıfakı and Erikli faults are major onshore shortening structures that appear to control much of the topographic relief in the Sinop Peninsula (Fig. 2A, B). The EeW striking Erikli Fault de-limits the low-relief topography of the peninsula and deﬁnes a prominent, straight mountain front, suggesting active deformation (Yildirim et al., 2011). The WNWeESE striking Balıfakı Fault runs through the central part of the peninsula and is located near the southern boundary of outcropping Neogene deposits (Fig. 2B). The EeW striking normal faults, immediately offshore from the Sinop Peninsula to its north and east (Fig. 2), deform Quaternary units (Özhan, 1989).

The basement rocks of the Sinop Peninsula consist of accreted late Cretaceous volcanic rocks, Maastrichtian-Paleocene and Eocene clastics and carbonates (Fig. 2B), which were amalgamated during the Alpidic orogeny during the Paleogene (Okay and Tüysüz, 1999). Miocene marine sediments exposed in the Sinop Peninsula include early-middle Miocene shallow-water limestones, sand-, silt-, and mudstones (Gedik and Korkmaz, 1984). They are gently folded into an open, EeW oriented syncline. These structures

Fig. 2. (A). Digital elevation model (25 m resolution) and morphotectonic segments of

the Sinop Peninsula. Onshore faults fromS¸enel (2002), offshore faults fromFinetti et al.

(1988)andÖzhan (1989). Red lines indicate paleoshorelines across the peninsula. Red

stars show OSL sampling locations. Capital letters and numbers in white boxes indicate

sample codes inTable 1. (B) Simpliﬁed geological map of the Sinop Peninsula (modiﬁed

afterS¸enel, 2002). Labels in boxes indicate tentative marine isotopic stages based on

our OSL measurements of terrace formation ages. (For interpretation of the references

to color in thisﬁgure legend, the reader is referred to the web version of this article.)

C. Yildirim et al. / Quaternary Science Reviews 81 (2013) 12e28 14

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indicate general regional, NeS directed, post-Tortonian shortening. The presence of Pliocene-Quaternary continental deposits super-seding the Miocene marine deposits indicates that the peninsula had already emerged above sea level in Pliocene time. Pleistocene and recent aeolian deposits form dune ridges atop most of the _Inceburun promontory at the NW corner of the peninsula (Fig. 2B). 3. Methodology

3.1. OSL age estimates

Optically Stimulated Luminescence (OSL) (Aitken, 1998) is a powerful dating tool for Quaternary sedimentary deposits, espe-cially when the age of deposition exceeds the time range of radiocarbon dating (i.e., 50 ka) or when suitable materials for alternative radiometric dating methods (e.g., U-series dating, cosmogenic nuclide dating, etc.) cannot be retrieved (Choi et al., 2008; Jacobs, 2008). An OSL age is obtained from the ratio of the paleodose (or equivalent dose) De in Gy accumulated in quartz over time during the burial of sediment due to the dose rate R (Gy/a) of the radiation environment:

OSL Age¼ De=R (1)

A rigorous OSL analysis requires complete bleaching (sufficient exposure to daylight) of quartz mineral grains during erosion and transportation prior to deposition. This condition is probably commonly met in the case of aeolian deposits, but it may not al-ways be achieved influvial and marine deposits due to the atten-uation effect of water on the intensity and spectral composition of daylight (Stokes, 1999; Jacobs, 2008; Fuchs and Lang, 2009). The stochastic nature offluvial and marine environments may give rise to differential and/or partial bleaching of the quartz grains, which would result in overestimated ages and/or variable age estimations. In addition to differential and/or partial bleaching, high concen-trations of water or carbonate content within the sediments can significantly decrease the dose rate (environmental radiation), resulting in overestimated ages.

We collected three samples to be dated with OSL from each of the sampled outcrops except for the 65_{2 m terrace in the} Boz-burun Promontory, from which we collected only one sample. 3.1.1. Sample preparation and OSL measurements

The sediment samples to be used for OSL dating were wet sieved to separate the grain size of 90e180

m

m. The separated grains were treatedﬁrst with 10% HCl and then H2O2for the removal of

car-bonates and organics, respectively. Then they were then etched with HF to remove the outer surface of the grains affected by alpha radiation, followed by HCl treatment once more. The samples were washed with distilled water and dried in an oven at 50C prior to OSL measurements. The purity of the quartz grains was tested by the absence of luminescence during infrared stimulation. All OSL measurements were performed with an automated Risø TL/OSL reader, model TL/OSL-DA-15, equipped with an internal 90Sr/90Y beta source (w0.1 Gy s1), blue light emitting diodes (LEDs) (470 nm,_{w40 mW cm}2) and IR LEDs (880 nm,_{w135 mW cm}2). Luminescence signals were detected using an EMI 9635QA photo-multiplier tubeﬁtted with 7.5 mm-thickness Hoya U-340 ﬁlters (Bøtter-Jensen, 1997). Sample preparation including chemical extraction of quartz from sediment material and OSL measure-ments were performed under subdued red light.

3.1.2. Equivalent dose (De) estimate

The equivalent dose (De) accumulated in quartz grains was

estimated using a single-aliquot regenerative-dose protocol

(OSL-SAR) based on a comparison of the natural OSL signal with regen-erative OSL signals produced by known laboratory doses (Murray and Mejdahl, 1999; Murray and Wintle, 2000). As an example, the radial plot of single grain Dedistributions of sample ICP-1A is

shown inFig. 3. A natural aliquot wasﬁrst preheated at 260_{C for}

10 s and then recorded with blue light stimulation at 125C for 40 s to obtain the natural OSL signal (Ln). To monitor and correct the

possible sensitivity change, a test dose was administered (10e20% of the natural dose) to the same aliquot prior to heating to 190C to release electrons from shallow traps. The test dose OSL signal was recorded to obtain (Tn). In the following three cycles, three

regen-eration doses were applied, following the same sequence of treat-ments as described in the ﬁrst cycle to obtain regenerated OSL signals (Li, i¼ 1,2,3). The same test dose was given again, heated to

190C, and the test dose OSL signal measured to give (Ti). In the

fourth cycle, bleached samples were measured to obtain the zero dose point (i¼ 0). Corrected OSL signals (Li/Ti, i¼ 1, 2, 3 and 0) were

used to construct the growth curve where the sensitivity-corrected natural signal (Ln/Tn) is interpolated onto the growth curve to

obtain De.Fig. 4A presents growth curves for the samples of ICP-2A,

B and C, where the corrected dose pointsﬁt well with the expo-nential function and the natural dose (open diamond) could be measured before saturation.

The aliquots were tested for sensitivity changes between the cycles. To do this, a known dose equal to thefirst regeneration dose was given to the samples during SAR application to obtain the regenerative dose points on the growth curves. The ratio of these two corrected regenerated OSL signals, namely the recycling ratio, is expected to be close to unity. InFig. 4BeD, the first two dose points indicate the same regenerative dose points, where solid circle shows the dose point from thefirst cycle and open circle from thefifth cycle of the same dose, which are rather close to each other on the growth curves.

As an internal reliability test the OSL measurements, a known dose (equal to theﬁrst regeneration dose) was given to the samples to recover the known dose. The recovered dose was found using the same SAR sequence as described above. The ratio of obtained dose to given dose should be close to unity, indicating that the dose measurement by OSL is reliable (Fig. 4E).

The dose rate was obtained from concentrations of the major radioactive isotopes of the U and Th series and of K in each sample (Olley et al., 1996). Radionuclide concentrations presented in

Table 1 were deﬁned using ICP-ES/ISP-MS analysis, by ACME Analytical Laboratories in Canada. The cosmic ray contribution to the dose rate was estimated for the altitude, latitude, longitude, and depth of each sample (Prescott and Hutton, 1988, 1994). The OSL ages, equivalent dose values and, dose rates obtained are presented inTable 1together with the number of aliquots evaluated for each sample.

Fig. 3. Radial plot of paleodose data from sample ICP- 1, where N represents the number of aliquots evaluated by OSL-SAR techniques. (RISOE Analyst program was used).

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3.2. Inner-edge altitude determination

An inner edge, or shoreline angle (line of intersection between the uplifted relict cliff and the relict wave-cut platform) of a marine terrace is the best geomorphic datum to calculate the relative elevation of past sea-level positions. The limited spatial extent of the terraces, locally dense bush vegetation, and slope degradation (e.g., Anderson et al., 1999) make it very difﬁcult to determine inner-edges at many locations from in situ measurements. In this study, we obtained our shoreline-angle elevation measurements based on (1) outcrop observations where stream valleys or roads cut the terraces, combined with (2) topographic proﬁles using GNSS differential Global Positioning System (dGPS) measurements.

We utilized a Matlab script that is integrated with ArcGIS to esti-mate best-_{ﬁt shoreline-angle elevations based on extrapolating the} slope of the relict cliff and the wave-cut platform along shore-normal proﬁles. The results typically yield 2-sigma errors in the range of 0.2e2 m. Errors are usually larger for the older, higher terraces.

3.3. Vertical displacement and uplift-rate calculation

Our chronologic interpretation of the marine terraces is based on the OSL dating results of marine sediments. Because this method tends to yield errors in age estimates on the order of 10%, comparing the OSL ages to the history of sea-level oscillations can

Table 1

Radiometric data and OSL ages. Terrace level Sample code Coordinates (UTM_36N) Depth (cm)

Dose (Gy) (N) Dose rate

(Gy/ka) Cosmic U (ppm) Th (ppm) K (%) Water Cont. (%) CaCO3 (%) Age (ka) ICP_A N_4662115 200 256 7 25 0.5 0.03 0.17 0.6 2.5 0.055 36 0 587 37

ICP-1 ICP_B E_660843 350 255 6 26 0.43 0.04 0.17 0.5 2 0.055 26 1 634 53

ICP_C 450 251 6 27 0.44 0.04 0.13 0.5 2 0.06 19 0 638 59

ICP_A N_4660348 50 34 1 8 0.53 0.04 0.23 0.6 2.1 0.04 16 0 75 5

ICP-2 ICP_B E_663209 100 36 1 8 0.61 0.04 0.2 0.6 2.85 0.07 22 0 67 5

ICP_C 150 33 1 8 0.45 0.04 0.19 0.5 2.1 0.055 21 0.6 74 6 SNP_A N_4656145 200 58 7 9 0.3 0.02 0.18 0.7 0.8 0.03 28 74.2 196 26 SNP-1 SNP_B E_679060 250 72 7 7 0.31 0.02 0.17 0.7 1.3 0.02 27 70 231 29 SNP_C 300 92 7 10 0.27 0.02 0.15 0.7 0.9 0.03 25 98.4 348 36 SNP-2 SNP_2 N_4656302 100 263 9 23 0.33 0.02 0.2 0.5 2 0.055 25 94.4 583 42 E_679539 125 5 AYN_A N_4645026 200 80 6 13 0.49 0.03 0.17 0.5 2.5 0.06 23 0 176 17

AYN-1 AYN_B E_639730 250 87 9 13 0.46 0.03 0.17 0.5 2.5 0.075 24 0 190 24

AYN_C 350 180 10 14 0.58 0.04 0.14 0.5 3.3 0.1 22 2.2 366 32

AYN_A N_4645302 300 178 8 15 0.47 0.04 0.16 0.5 2.6 0.05 32 1 405 36

AYN-2 AYN_B E_645612 150 167 11 12 0.53 0.04 0.19 0.5 2.3 0.07 32 0 374 36

AYN_C 250 230 9 25 0.51 0.04 0.16 0.5 2.6 0.055 27 0.08 538 45

ICP-2 samples are from a dune at the top of the Inceburun promontory.

Fig. 4. (A) Dose response curves constructed using corrected OSL signals obtained for samples ICP-2A, ICP-2B and ICP-2C. Curves and Dedose are similar for these three samples. (B),

(C), (D) Dose response curves constructed using corrected OSL signals obtained for the samples AYN-2C, ICP-1A and SNP-2. Open diamond indicates the natural dose point on the growth curve; the interpolation of this point over growth curve to horizontal axis gives the OSL SAR dose, where they are 211.3 Gy, 279.3 Gy and 284.5 Gy, respectively. Open and solid circles show the repeated dose points. (E) Recycling ratio values for ICP samples together with error bars.

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provide more precise age estimates. The sea-level highstands that form marine terraces are associated with interglacial stages and interstadials, which are correlated with odd-numbered marine isotopic substages (MIS) (Bloom et al., 1974;Pirazzoli et al., 1993; Muhs et al., 2002; Pedoja et al., 2011). Where the uplift rate is high enough, terraces related to short-lived interstadial sea levels may be preserved onshore (e.g., Lajoie, 1986). To determine the most likely highstands that correspond to the formation of the marine terraces, we tentatively correlated the OSL ages of the marine terraces to the chronologically closest marine isotopic stages using the global eustatic sea-level curve ofBintanja et al. (2005)and Siddall et al. (2006), which cover the time range of our OSL ages and are widely accepted. We use these curves instead of the Black Sea sea-level curve (Shmuratko, 2001), because the sill depth of the _Istanbul Strait (32 m) is topographically lower than sea-level highstands in the Pleistocene. In fact, connectivity be-tween the Black Sea and the Mediterranean (e.g.,Zubakov, 1988) can be inferred from a recently published stacked speleothem

d

18O record from Sofular Cave (250 km west of the Sinop Peninsula), which indicates the possibility of twelve separate marine trans-gression events from the Mediterranean to the Black Sea via the shallow _Istanbul Strait during the last 670 ka (Badertscher et al., 2011).

In general, the net vertical displacement of the marine terraces is calculated by using the difference between the present-day terrace elevation (E) and the paleoelevation (e) of each shoreline.

To calculate the mean uplift rate (U) of each terrace, the net vertical displacement is divided by the MIS age (A) of the marine terrace (Lajoie, 1986):

U ¼ ðE eÞ=A (2)

Nevertheless, differences among the different paleo sea-level curves with regard to the timing, height, or depth of the MIS highstands makes it difﬁcult to deﬁne a precise altitude of the paleoshorelines to calculate net vertical displacements (Caputo, 2007). Therefore, we followPedoja et al. (2011)and calculate the apparent vertical displacement rates without eustatic corrections. The apparent displacement rate is calculated by only using the present-day terrace elevation (E) and MIS age (A) of the marine terrace (Lajoie, 1986):

U ¼ E=A (3)

Errors reported on the apparent displacement rates include er-rors in the MIS ages and in inner edge altitude measurements (GNSS) of the terraces. For comparison, we also calculate net ver-tical displacement with corrections for eustatic sea-level based on the eustatic sea level curves ofShmuratko (2001), Bintanja et al. (2005)andSiddall et al. (2006), and present them in a compara-tive plot that illustrates the differences in uplift rate estimated with each curve.

Fig. 5. (A) Image of the sampling location and uplifted shorelines at the _Inceburun Promontory. White-colored dashed lines and numbers indicate proﬁle line and corresponding

numbers, respectively. Two-headed arrows indicate viewing direction of photographs inFig. 4AeC. (B) Topographic proﬁles and inner edge elevations of the sampled marine

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4. Results

4.1. Geomorphology of marine terraces

Severalﬂights of marine terraces are eroded into the Upper Cretaceous to Miocene rocks within the different morphotectonic segments of the Sinop Peninsula. The main morphotectonic units in the Sinop Peninsula include (i) the _Inceburun Promontory; (ii) the Bozburun Promontory; and (iii) the AyancıkeGerze Zone (Fig. 2A). 4.1.1. _Inceburun Promontory

The _Inceburun Promontory is characterized by low-relief topography between the Black Sea and the Balıfakı Fault (Fig. 2A). Inactive dune_{ﬁelds cover extensive areas of the terrace surfaces at} the top of the promontory, especially in the north (Fig. 2A). The coast comprises rocky shores with 10 to 40-m-high cliffs and es-tuaries and rias similar to the Karasu coastal plain (Fig. 2A). The sequence on the promontory includes two levels (T1 and T2) of marine terraces with inner edge elevations between 5 1 and 12 0.7 m that have been sculpted into upper Cretaceous basalts in the northernmost part of the promontory (Figs. 2B, 5 and 6AeC). The wave-cut notch of the 5 1 m paleoshoreline can be observed in the modern cliff at the northernmost part of the peninsula (Fig. 6C). We also observed wave-cut notches and a shingled beach over a shore platform associated with the same 5 1 m elevation paleoshoreline in an isolated bay of the western part of the peninsula (Fig. 6B). The outer edge of the abrasion platform and deposits of the 12 0.7 m terrace are well exposed in the north-western part of the peninsula (Fig. 6A). The top surface of this terrace is veryﬂat at an elevation of 17 0.4 m; the surface is covered a thick soil horizon, indicating long-acting pedogenic processes.

4.1.2. Bozburun Promontory

The Bozburun Promontory is a small peninsula immediately east of the town of Sinop (Fig. 2A). The promontory has relatively high relief with rocky shores, 20- to 150-m-high sea cliffs, and a smooth, southward tilted top surface that reaches up to 250 m elevation. The Bozburun Promontory is connected with the _Inceburun Promontory via a 3-km-long isthmus (Fig. 2A), which includes a 2-to 5-m-thick layer of marine sediments rich in shell fragments. The marine terrace sequence on the Bozburun Promontory consists of four levels with inner edges at elevations of 4_{0.2, 7 0.5 m,} 17 0.4 m, 34 2 m and 65 2 m (Fig. 6DeF and 7A, B) that have been eroded into the upper Cretaceous basalts. The upper parts of the paleocliffs are still distinguishable, but thick colluvial aprons cover the cliff bases, the inner parts of the terrace treads and the shoreline angles. Nevertheless, the outer edges of the terraces are still preserved. The wave-cut notch of the 7 0.5 m level and an associated platform are eroded into the modern cliff (Fig. 6D), suggesting recent relative sea-level fall. The top surface of the 17 0.4 m level reaches 25 m elevation as a result of colluvium derived from the 65 2 m terrace. We couldn’t ﬁnd any outcrop associated with marine deposits at this terrace level, but its outer edge is exposed at 15 m elevation. The top surface of the 34 2 m terrace level is the largest among the terraces. The elevation of this level reaches 40 m toward the inland, while its outer edge is 28 m

on the modern cliff. The highest level at 65_{2 m is associated with} marine sediments up to 2 m in thickness and an outer edge elevation of 64 m.

4.1.3. Ayancık-Gerze zone

The Ayancık-Gerze zone is transitional between the _Inceburun Promontory and the Sinop Range, between the Erikli and Balıfakı faults (Fig. 2A and B). This zone has higher topography and relief compared to the _Inceburun Promontory (Fig. 2A). The largest ma-rine terraces are found east of the town of Ayancık and north of Gerze, along the western and eastern shores of the hanging-wall block of the Balıfakı Fault (Fig. 6G and H). Here, the coast is pre-dominantly rocky, with 10- to 70-m-high sea cliffs. Immediately east of Ayancık, the sequence consists two levels of marine terraces with shoreline angles at elevations between 21 0.7 m and 67 1.4 m (Fig. 8A and B) that were carved into clastic and car-bonate rocks of Maastrichtian-Paleocene age (Fig. 2B). Marine sediments 2- to 10-m thick locally cover abrasion platforms of the terraces. The most prominent terrace level is 67 1.4 m high, covering an area 0.5-m to 1.5-km wide and 7- to 8-km long, which is delimited to the east by the Balıfakı Fault (Fig. 2A). The top surface of the terrace reaches up to 75 m, where a colluvial apron covers the inner edge of the terrace. The elevation of the outer edge associated with this level is 45 m at the modern cliff. The lower terrace level has relatively small surface west of the higher terrace level. The elevations of the top and outer edge of the terraces are 25 m and 15 m, respectively. The thicknesses of the marine deposits vary between 3 and 5 m.

4.2. Stratigraphy of the marine terrace deposits

Although the marine terraces constitute abrasion platforms, marine sediments cover virtually all of them. In the following sections we describe these deposits, including those which we sampled for OSL dating.

4.2.1. _Inceburun Promontory

The _Inceburun stratigraphic section is exposed in the north-ernmost part of the Sinop Peninsula. This section comprises a 6-m-thick sedimentary package without fossils, organized in upward-fining sandstone sequences (Figs. 6A and 9A and B). The base of the section consists of a 1-m-thick, fine-grained sandstone with horizontal to sub-horizontal lamination. Up-section, these layers alternate with 1- to 3-cm-thick medium- to coarse-grained sand-stone. These strata are overlain by up to 50-cm-thick channelized, medium-grained, horizontally laminated sandstones with a pro-nounced down-cutting erosive base, grading up-section into an alternation of medium- and coarse-grained, horizontally-lami-nated, 1- to 5-cm-thick sandstone layers. Toward the top is a chaotic cross-bedded horizon, associated with a short-wavelength, dis-harmonic, and discontinuously laminated sandstone (Fig. 9A and B). Immediately above this horizon, 0.2-m-thick, steeply cross-bedded strata are preserved with a convex-upward geometry on the crest and concave-upward in the swale. Finally, in the upper-most part of the section, weathered and mottled _{fine-grained} sandstone prevails.

Fig. 6. (A) View of the ICP1 sampling location and exposed part of the abrasion platform of the terrace in the northern part of the _Inceburun Promontory (seeFig. 3for point of

view). (B) View of the 4- to 5-m marine terrace inside a small bay. Pink dashed line indicates wave-cut notch of the paleoshoreline. (C) View of the 4- to 5-m wave-cut notch in the

northernmost part of the _Inceburun Promontory (seeFig. 3for point of view). (D) View of the 7 0.5 m and 17 0.4 m terraces in the northernmost part of the Bozburun

Promontory. (E) Close-up view of the wave-cut notch of the 7 0.5 m terrace. (F) View of the SNP1 and 2 sampling locations and exposed part of the abrasion platform of the terrace

north of the Bozburun Promontory (seeFig. 5for viewing direction). (G) and (H) view of terraces in the Ayancık-Gerze zone (seeFig. 6for viewing direction). Capital letters and

numbers in the upper boxes indicate marine isotopic stages and numbers in the white boxes indicate shoreline-angle elevations of the sampled terraces. (For interpretation of the

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The sedimentary facies association suggests deposition in a shallow-water marine environment within a general trend toward increased water depth through time. The channelized sandstone at the base of the section suggests tidal inﬂuence, while the chaotic horizon is interpreted to represent slumping (Fig. 9C), and hence slope instability. The cross-bedding geometry preserved on top of the slumped horizon suggests that the slumped unit could repre-sent swaley cross stratiﬁcation (SCS), indicating deposition above the storm-wave base (e.g., Dumas and Arnott, 2006). Massive mudstone typical of an offshore depositional environment, how-ever, is absent. Therefore, despite the overall deepening trend of the depositional environment, the section probably terminates above of the shorefaceeoffshore transition.

4.2.2. Bozburun Promontory

Terraces in the Bozburun Promontory are capped by marine sediments from middle to late Pleistocene age.Erinç and _Inandık (1955), _Inandık (1955), Akkan (1975), and Karabıyıkoglu (1984)

studied the late Pleistocene coastal deposits in detail. The sec-tions are not well exposed today due to intense urbanization; therefore, we refer to the previously published descriptions. The thickness of the sedimentary succession is about 40 m along the isthmus. The succession consists of a 4- to 5-m-thick cross-stratiﬁed coquina sandstone at the base overlain by a 18- to

20-m-thick high-angle, cross-stratified sandstone, followed by a 12-to 15-m-thick massive muds12-tone, and_{finally a 4- to 5-m-thick} high-angle cross-stratified sandstone at the top (_Inandık, 1955;

Karabıyıkoglu, 1984). Karabıyıkoglu (1984)interpreted the sand-stones as transgressive lag deposits in a foreshore/backshore environment associated with coastal dunes. The contact with Pre-Quaternary deposits is not exposed.

In the rocky coast part of the promontory, the 6- to 7-m-thick low-angle cross-stratiﬁed shelly sandstone overlies 1- to 2-m-thick gravels (Erinç and _Inandık, 1955; _Inandık, 1955; Karabıyıkoglu, 1984), which in some places are deposited directly over elevated abrasion platforms in the basalts. Together, the facies correspond to shoreline deposits.

4.2.3. Ayanc_{ık-Gerze zone}

The Ayancık stratigraphic section is characterized by a 3-m-thick sedimentary sequence made up of packages of 20- to 50-cm-thick fine- to medium-grained, amalgamated sandstone, alter-nating with 1- to 5-cm-thick coarser strata with isolated pebbles (Fig. 10A and B). The 20- to 50-cm-thick sandstone layers have a complex cross-stratification pattern, with sharp surfaces at the base and top of each package, horizontal lamination, cross lamination, and parallel lamination toward the top, with a convex-upward and a concave-downward stratification still preserved (Fig. 10C).

Fig. 7. (A) Image of the sampling location and uplifted shorelines in the Bozburun Promontory. White dashed lines and numbers indicate line of proﬁle and corresponding numbers,

respectively. Two-headed arrows indicate viewing directions in photos ofFig. 4. (B) Topographic proﬁles and inner edge elevations of the sampled marine terraces.

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Similar to the Bozburun section, the facies association of the Ayancık sedimentary sequence suggests deposition in a shallow-water marine system. We interpret the cross stratiﬁed sandstone to represent swaley cross-strata (SCS), thus reﬂecting deposition above the storm-wave base. The relatively good preservation of the SCS and the amalgamation suggest a relatively high sedi-mentation rate.

4.2.4. Paleobathymetry estimations

Sedimentary facies and structures do not unequivocally reﬂect a speciﬁc water depth, but they represent the interaction between the sediment characteristics (i.e., grain size, shape, density, composition etc.) and the hydrodynamics of the water system (e.g.,

Bridge and Demicco, 2008; Immenhauser, 2009). Therefore, assigning paleodepths based on sedimentary facies associations alone is problematic and may introduce signiﬁcant uncertainty.

The stratigraphic section of the _Inceburun promontory reflects a deepening-upward trend characterized by a fining-upward sequence, including sedimentary structures typical of tidal chan-nels and swaley cross-stratification (SCS). The water depth of a tidal channel _{floor may range between 5 and 30 m (}Immenhauser, 2009), while the SCS, which is thought to reflect a combination of oscillatory and unidirectionalflow, should represent deposition in proximity of the fair-weather wave base (e.g.,Dumas and Arnott, 2006). Sedimentation at that depth is consistent with the lack of massive mudstone and the presence of isolated pebbles. To our knowledge, however, there are no estimations of the average depth of the fair-weather and storm-wave bases in the present-day Black Sea. A global compilation indicates that fairweather wave-base depths for shallow siliciclastic seas are 10 5 m (e.g.,

Immenhauser, 2009). For the _Inceburun section, this would represent a minimum estimate, because the section shows a deepening trend toward the offshore transition region. In addition, this sedimentary unit is located in proximity to the respective paleocliffs of the _Inceburun Peninsula, suggesting that the elevation of our sampling site (10 2 m) is very close to the elevation of the respective paleoshoreline (12 0.7 m) obtained from dGPS mea-surements. Together, these observations imply a water depth of 1e 5 m, which is slightly lower than the estimation (10 5 m) compiled by Immenhauser (2009). Similarly, the well-preserved SCS (Fig. 10) in the Ayancık sedimentary sequence suggests depo-sition in proximity to the fairweather wave base (e.g.,Dumas and Arnott, 2006). In addition, our GPS survey documents a 13- to 20-m difference between the elevation of our sa20-mpling site (50 2 m) and the elevation of the paleoshoreline (67 1.4 m). This difference is still similar to the global fairweather wave-base estimate (10 5 m;Immenhauser, 2009).

Based on these observations, and considering the intrinsic dif-ﬁculty related to water-depth estimates, we consider a paleodepth of 5e10 m for both stratigraphic sections. This conservative approach, which is consistent with the different estimations, avoids introducing further uncertainties in the uplift rates among the different sites.

4.3. OSL ages and assigned MIS stages of the marine terraces

Quartz obtained from all sediment samples was bright in luminescence, and sensitivity changes were corrected successfully using the response to a test dose. The OSL samples that we collected from marine terrace deposits (calculated using Eq.(1)) yield ages

Fig. 8. Geomorphic map of the marine terraces in the eastern part of the Ayancık-Gerze Zone. White dashed lines indicate line of proﬁle with corresponding numbers. Two-headed

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that range from 176 17 ka to 638 59 ka and fall within several clusters (Table 1). These relatively old OSL ages result from low environmental radiation as well as high equivalent dose values, and they appear to be reliable based on growth curves and dose re-covery tests.

From an outcrop close to the inner edge of the 12 0.7 m terrace of the _Inceburun Promontory (level ICP-1), the three samples yielded ages of 587 37 ka, 634 53 ka, and 638 59 ka (Table 1). OSL ages from the 34 2 m terrace at the Bozburun Promontory (level SNP-1) are 196 26 ka, 231 29 ka, and 348 36 ka, the latter of which was from the outer edge of the terrace. For the 65_{2 m terrace at the Bozburun Promontory (level SNP-2), our} sample from the outer terrace edge yielded an age of 583 42 ka (Table 1). In the Ayancık-Gerze Zone, the OSL ages from the outer edge of the 21 0.7 m terrace (level AYN1) are 176 17 ka, 190 24 ka, and 366 32 ka (Table 1), while ages from the inner edge of the 67 1.4 m terrace (level AYN2) are 374 36 ka, 405 36 ka, and 538 45 ka (Table 1).

Different ages obtained from a single continuous section without any unconformities might imply differential and/or partial bleaching of the grains. In those cases, the ages of the samples

obtained from the minimum equivalent dose (De) value might be more trustworthy due to more complete bleaching. Of course, these samples could still be partially bleached and we may overestimate the original age of the terrace. Therefore, we consider the age of the deposit to represent the maximum age of the wave-cut platform, and because its formation can precede or be coeval with deposition, the uplift rates that we obtain represent minimum rates.

Based on this argument, we assigned each sample location to a corresponding sea-level highstand by using the age of the sample having the minimum equivalent dose (De) for each continuous section. Accordingly, we assign the 12 0.7 m terrace in the _Inceburun Promontory (ICP-1) to MIS 15a-15c (ca 570 ka), and the 34 2 m (SNP-1) and 65 2 m (SNP-2) terraces along the Boz-burun Promontory to MIS 7a (ca 190 ka) and MIS 15a-15c (ca 570 ka), respectively. We also identiﬁed a terrace level at 17 0.4 m and a wave-notch level at 7 0.5 m in the Bozburun Promontory (Fig. 8,“SNP-0” inTable 2), which may correspond to MIS 5e (ca 125 ka) because of its position immediately below the MIS 7a terrace level. We assign the marine terrace levels at 21 0.7 m (AYN-1) and 67 1.4 m (AYN-2) east of the town of Ayancık to MIS 7a (ca 190 ka) and MIS 11 (ca 400 ka), respectively

Fig. 9. (A) Stratigraphic section of the ICP1 sampling location in the northern part of the _Inceburun Promontory. Numbers in the boxes indicate our OSL ages from the section. (B) General view of the sampled outcrop. Holes show our sampling points. (C) Close-up view of the slump structures in the section that might be a result of seismic shaking.

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(Table 2,Fig. 11). Although in some cases we obtained ages for in-dividual terrace levels that could correspond to different odd numbered MIS substages (for example, SNP-1 and AYN-1, AYN-2 and ICP-1), we did not observe any sedimentary cannibalism that might indicate a marine transgression, erosion, and deposition related to several MIS highstands.

The only other geochronologicalﬁnding relevant to our terrace ages is the biostratigraphy related to marine shell fragments that we identiﬁed along the isthmus of the Bozburun Promontory.Erinç and _Inandık (1955)described these marine shells as a characteristic fossil assemblage of the Karangatian Transgression in the Black Sea that corresponds to the MIS 5e highstand (e.g.,Tchepalyga, 1997). The isthmus where they found the characteristic fossil assemblage is one terrace level lower than the terrace we relate to MIS 7a from our OSL ages, which helps to validate the accuracy of our OSL ages and corresponding MIS assignments, at least in the Bozburun Promontory.

In addition to samples from marine terraces, we sampled a paleodune ridge overlying the marine deposits at the northernmost part of the peninsula (Level ICP-2 inTable 1, andFig. 2A). The OSL samples ICP 4, 5 and 6 yielded ages of 75 5 ka, 67 5 ka, and 74 6 ka, which we associate with the sea-level low stand at the transition from MIS 5 to 4 (Table 1).

4.4. Vertical displacements and mean-uplift rates of the marine terraces

The vertical displacements of the terraces that we report and discuss in this study are the apparent vertical displacements of the terraces. The net vertical displacements of the terraces, which include a correction according to the sea-level curves ofShmuratko (2001), Bintanja et al. (2005),Siddall et al. (2006), are also given in

Table 2andFig. 12to show the differences between the two ap-proaches. The apparent vertical displacements range from 12 m to 67 m (Table 2). The lowest and highest vertical displacements are obtained from the 12 0.7 m (MIS 15) and 67 1.4 m (MIS 11) terraces in the _Inceburun Promontory and Ayancık-Gerze Zone, respectively. The apparent vertical displacement of the MIS 15 terrace level of the Bozburun Promontory is 53 m greater than that of the terrace of the same age in the _Inceburun Promontory mor-photectonic segment (Table 2).

The lowest mean apparent vertical displacement rate is 0.02 mm/a (since ca 570 ka), obtained from the _Inceburun Prom-ontory in the northernmost part of the Sinop Peninsula (Table 2,

Fig. 13). At the Bozburun Promontory, the uplift rate from 570 to 190 ka was 0.08 mm/a, then increased to 0.26 mm/a from 190 to 125 ka, and slowed to 0.14 mm/a from 125 ka to today (Table 2,

Fig. 10. (A) Stratigraphic section of the AYN2 sampling location in the Ayancık-Gerze Zone. Numbers in the boxes indicate our OSL ages from the section. (B) General view of the sampled outcrop. (C) Close-up view of the swaley cross stratiﬁcation in the section.

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Figs. 13and14). In the Ayancık-Gerze morphotectonic segment, the uplift rate started off at 0.22 mm/a from 400 to 190 ka, and then decreased to 0.11 mm/a since ca 190 ka (Table 2,Figs. 13and14). This slowing of the uplift rate might be the reason why we observe the MIS 7a level at lower elevation (21 0.4 m) in the Ayancık-Gerze Zone with respect to its counterpart (34 2.1 m) in the Bozburun Promontory.

As depicted inFig. 12, the application of an eustatic correction to estimate vertical displacement rates from marine terraces might include major biases, particularly when estimates are made from terraces of different ages. This is particularly the case in areas where the uplift rate is relatively low (<0.2 mm/yr), such as the Turkish coast of the Black Sea. Sea-level curves have been determined using variable methods at different locations, where the local effects of dynamic topography and other tectonic forces might influence the relative amplitude of sea-level oscillations; these effects are dif fi-cult to quantify (e.g.,Pedoja et al., 2011; Rowley et al., 2013). In fact, the vertical displacement rates estimated using the sea-level curve ofShmuratko (2001), which is specific to the Black Sea, are most similar to the apparent displacement rates, thus suggesting no major local bias. However, the general trend of northward-decreasing vertical displacement rates across the Sinop Peninsula is independent of the use of a eustatic correction (Fig. 13). 5. Discussion

5.1. Non-uniform uplift of marine terraces

The spatiotemporal variations in apparent uplift rates indicated by the ages of the marine terraces can provide insights into the deformation history of the Sinop Peninsula and the outer parts of the orogenic wedge bordering the northern sector of the Central Anatolia Plateau. The terraces along the _Inceburun Promontory at the NW tip of the Sinop Peninsula yield the lowest uplift rates among our data (0.02 mm/a since 570 ka) (Fig. 14). Interestingly, between ca 570 and 190 ka, the Bozburun Promontory at the NE tip records an uplift rate of 0.08 mm/a, which is slightly faster but still similar to the long-term uplift rate at the _Inceburun Prom-ontory (Table 2,Fig. 14). Nevertheless, the uplift rate of the Boz-burun Promontory increased to 0.26 mm/a between ca 190 and 125 ka, which is three times higher than the previous uplift rate, then slowed to 0.14 mm/a since 125 ka (Table 2, Fig. 14). Comparing the two sites, we interpret the pre-190 ka uplift history to indicate that both regions were slowly, but nonetheless differ-entially uplifted above an offshore thrust fault that is below and to the north of both areas. After 190 ka, the faster uplift of the Boz-burun Promontory likely indicates higher activity along the offshore structures that lie immediately to the south and north (Fig. 2), which Özhan (1989) interpreted to be normal faults bounding a horst. Compared to the _Inceburun and Bozburun promontories, the Ayancık-Gerze zone, which is in the hanging-wall block of the Balıfakı Fault, records relatively fast uplift with a rate of 0.22 mm/a from 400 to 190 ka, but slowed to 0.11 mm/a since 190 ka (Fig. 14). We relate these uplift rates to movement along the Balıfakı Fault, as it is the only mapped structure that could potentially accommodate differential uplift between the Ayancık-Gerze zone (at the southern end of the Sinop Peninsula) and the promontories to the north.

We thus conclude that shortening deformation has predomi-nated during the Quaternary in the Sinop region. In this context, we interpret the spatially limited occurrence of normal faulting at the northern and southern boundaries of the Bozburun Promontory as a local phenomenon (see below). Fast uplift of the Bozburun Promontory between ca 190 and 125 ka could be associated with a transient phase of normal fault accommodation in the region,

T able 2 T entati v e marine iso topic stages, to ta l displacements and uplift ra te s o f the te rr aces. Morphotectonic segment Terrace level Assigned MIS stage MIS age (ka) Inner edge elevation (m) Sea-level highstand a elevation (m) Sea-level highstand b elevation (m) Sea-level highstand c elevation (m) Total displacement (m) a Total displacement (m) b Total displacement (m) c Uplift rate a (mm/a) Uplift rate b (mm/a) Uplift rate c (mm/a) Uplift rate d (mm/a) Inceburun Promontory ICP_1 15 570 10 13 0.2 15 1 8 2 þ 1 21 28 1.0 21 2.0 12 1.0 0.05 0.01 0.04 0.01 0.02 0.003 0.02 0.003 Bozburun SNP_0 5e 125 51 7 0.4 þ 1 1 þ 5 2 þ 2 11 16 1.1 12 2.0 15 1.1 0.13 0.03 0.10 0.03 0.12 0.03 0.14 0.03 Promontory SNP_1 7a 195 53 4 2.1 20 4 11 40 41 54 4.5 45 4.5 34 2.3 0.28 0.05 0.23 0.04 0.17 0.03 0.18 0.03 SNP_2 15 570 10 65 2 15 1 6 3 þ 1 18 0 2.2 71 3.6 64 2.0 0.14 0.02 0.12 0.02 0.11 0.02 0.11 0.02 Ayancik-Gerze AYN_1 7a 195 52 1 0.7 20 4 11 4 þ 0 11 41 4.1 32 4.1 21 1.2 0.21 0.04 0.16 0.03 0.11 0.02 0.11 0.02 Zone AYN_2 11 400 56 7 1.4 1 1 2 1 1 11 68 1.7 69 1.7 66 1.7 0.17 0.02 0.17 0.02 0.17 0.02 0.17 0.02 a According to Bintanja et al. (2005) . b According to Siddall et al. (2006) . c According to Shmuratko (2001) . d Apparent uplift rate (e.g., Lajoie, 1986; Pedoja et al., 2011 ).

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which does not appear to be representative of the longer-term rate and style of deformation.

5.2. Implications for deformation patterns along the Pontide Belt The Pontide ranges are a wedge-shaped orogenic belt spanning nearly the entire southern (Turkish) coast of the Black Sea. The topography of the Pontides changes dramatically from west to east. The coastal zone mainly consists of erosional rocky shores where

severalﬂights of marine terraces are preserved (Erinç and _Inandık, 1955; _Inandık, 1955; Akkan, 1975; Keskin et al., 2011). Nevertheless, reliable marine terrace data, including absolute dating constraints and paleoshoreline elevations that are comparable with our data, are only available from the eastern Pontides. Seven levels of marine terraces uplifted between 3 and 260 m were mapped in the Trabzon area (Keskin et al., 2011), which is 400 km east of the Sinop Peninsula. These terraces were dated using ESR, and associated

Fig. 11. Global sea level curves (lower panel) for the last 700 ka, heights, and OSL-derived ages of the terraces (upper panel) across the Sinop Peninsula. Solid gray line indicates modern sea level, dashed line indicates sill depth of the _Istanbul Strait, which connects the Black Sea to the Mediterranean. Vertical scales of the sea-level curve and terrace heights are different. Numbers and letters on the eustatic sea-level curve (blue, black, and red lines) indicate marine isotopic stages and their sub-stages. (For interpretation of the

ref-erences to color in thisﬁgure legend, the reader is referred to the web version of this article.)

Fig. 12. Mean vertical displacement rates for marine terraces of the Sinop Peninsula. Rates have been calculated at sites with age constraints from OSL dating using eustatic corrections from three different sea-level curves; apparent rates include no eustatic correction (see text for details). Note that displacement rates estimated using the sea

level curve ofShmuratko (2001), which is speciﬁc for the Black Sea, are very similar to

the apparent estimates. The global sea-level curves ofBintanja et al. (2005)andSiddall

et al. (2006)and tend to include large biases because they include larger amplitudes of

the sea-level oscillations. In general, all estimates show a northward-decreasing pattern, which we associated with activity of local faults.

Fig. 13. Topography, faults (Özhan, 1989; S¸enel, 2002), morphotectonic segments, and uplift rates derived from marine terraces of the Sinop Peninsula. Red lines indicate paleoshorelines across the peninsula. Red stars denote OSL sampling locations.

Numbers in white boxes indicate uplift rates reported inTable 1. (For interpretation of

the references to color in thisﬁgure legend, the reader is referred to the web version of

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with MIS 11, 9 and 5e, resulting in mean uplift rates that decrease through time from 0.17 0.03 mm/a, to 0.10 0.02 mm/a, and ﬁnally 0.07 0.05 mm/a. The gradually decreasing uplift rates are lower than those that we obtained over a shorter timescale at the Bozburun Promontory and from the Ayancık-Gerze Zone (Table 1). These results are somewhat surprising, as one might expect higher uplift rates in the Eastern Pontides, considering their higher relief. Nonetheless, the slower uplift rates from the Eastern Pontides are consistent with close to neutral NAF-normal slip rates suggested from an analysis of regional GPS data (Reilinger et al., 2006). The relatively high uplift rates that we obtained from the Central Pon-tides might be associated with their position at the apex of the large restraining bend in the NAF, where contractional strain is likely accumulating (Yildirim et al., 2011).

5.3. Implications for deformation patterns in the Central Pontides

The patterns of uplifted marine terraces at the Sinop Peninsula indicate active internal deformation and upper crustal shortening within the Central Pontide orogenic wedge (Fig. 15). This defor-mation, however, is not only restricted to the Sinop Peninsula, but extends across the entire array of mountain ranges and intermon-tane basins between the NAF and the Black Sea (Yildirim et al., 2011, 2013).

The seismic reflection profiles from offshore Sinop, including areas on the continental shelf and the abyssal plain of the Black Sea, display different fault kinematics (Finetti et al., 1988; Özhan, 1989; Aydın et al., 1995). The abyssal plain has been deformed by imbri-cated thrusts, with younging of the deformation toward the hin-terland (Finetti et al., 1988). On the hanging-wall block of the youngest thrust of these imbricated structures, a set of normal faults deforms Quaternary units of the continental shelf to the north and east of the Sinop Peninsula (Özhan, 1989). In contrast to these normal faults, the Balıfakı thrust records shortening in the southernmost part of the Sinop Peninsula. Although normal faults deform the continental shelf, the northern margin of the peninsula is contractional overall, and localized normal faulting might be a result of localflexural bending-moment stresses in the hanging-wall block of the thrust faults that deform the abyssal plain (Fig. 15); alternatively, they might be related to gravitational collapse at the edge of the continental shelf.

The character of on- and offshore shortening across the Pontides since the late Miocene indicates out-of-sequence thrusting (e.g.,

Yildirim et al., 2011), where either a progressive forward or break-back sequence of thrusting implies that upper crustal shortening mainly takes place on faults in the interior part of the orogenic wedge. However, the active deformation of structures within and to the north of the Sinop Peninsula may indicate that since the Pleistocene, deformation in the orogenic wedge has been taking

Fig. 14. Uplift rates through time derived from the elevated Sinop marine terraces and MIS-stage assignments based on our OSL dating. Numbers on top indicate MIS stages.

Fig. 15. Conceptual diagram of the orogenic wedge of the Pontide mountains showing principal north-vergent thrust faults and Quaternary mean-uplift rates in the different fault blocks, emphasizing a really extensive shortening and differential uplift in different sectors of the orogenic wedge. Numbers above the colored lines indicate time over which the

uplift rate is integrated. The uplift rate in the Gökırmak Basin indicates mean uplift rate of the basin derived from ﬂuvial strath terraces (fromYildirim et al., 2013), not the faults.

Note the area of localized thin-skinned normal faulting in the hanging wall of the thrust fault of the northern Sinop Peninsula. AGZ: Ayancık-Gerze Zone, BP: Bozburun Promontory, _IP: _Inceburun Promontory, NAF: North Anatolian Fault, EkF: Ekinveren Fault, EF: Erikli Fault, BF:Balıfakı Fault.

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place over a broader region, extending between the interior of the wedge in the Kastamonu region to the south and offshore Sinop to the north.

6. Conclusions

Our sedimentologic and geomorphic observations, the compi-lation of regional structural data, and OSL dating of sediments capping marine terraces along the southern coast of the Black Sea supports the notion of overall shortening across the Sinop Penin-sula as an integral part of the activity of the Pontide orogenic wedge. Movements along these faults have accommodated differ-ential uplift over the past 570 ka. Uplift rates that range from ca 0.02e0.26 mm/a highlight the spatially non-uniform and temporally variable nature of regional uplift. The northern (_Inceburun and Bozburun) promontories experience similarly slow uplift rates of 0.02 and 0.08 mm/a from 570 to 190 ka, probably related to a slow-moving fault north of both sites. After 190 ka, the Bozburun Promontory alone experienced a phase of the fast (0.26 mm/a) uplift, which was likely linked to nearby offshore normal faults. In contrast, the Ayancık-Gerze zone in the southern part of the Sinop Peninsula experienced moderate uplift rates of 0.22 mm/a from 400 to 190 ka and 0.11 mm/a since 190 ka, which indicates tectonic activity along the Balıfakı thrust fault. Despite the late Miocene pattern of back-stepping deformation within the orogenic wedge and Quaternary reactivation of thrust faults farther inland in the Gökırmak Basin (e.g.,Yildirim et al., 2011, 2013), active shortening across the Sinop Peninsula recorded since the middle Pleistocene documents that the active deformation zone has expanded outward along the northern margin of the Central Anatolian Plateau.

Acknowledgments

This study is part of the Vertical Anatolian Movements Project (VAMP), funded by the TOPO-EUROPE initiative of the European Science Foundation, including contributions by the German Science Foundation (DFG) to MS (grants STR373/20-1 and 25-1). CY was supported by Deutsche Forschungsgemeinschaft (DFG: EC 138/5-1), Potsdam University, Deutsches GeoForschungsZentrum Pots-dam and TUBITAK 107Y333; HE was supported by Deutsche For-schungsgemeinschaft (DFG: EC 138/5-1). TS was supported by the Leibniz Center for Surface Processes and Climate Studies at the University of Potsdam (DFG: STR373/20-1) and the Alexander von Humboldt Foundation.

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